Supercontinent cycle

Wilson cycle

Simplified sketch of the western part of Pangaea

The supercontinent cycle describes the quasi-periodic aggregation and dispersal of Earth's continental crust. There are varying opinions as to whether the amount of continental crust is increasing, decreasing, or staying about the same, but it is agreed that the Earth's crust is constantly being reconfigured. One complete Supercontinent cycle is said to take 300 to 500 million years.

Continental collision makes fewer and larger continents while rifting makes more and smaller continents. The last supercontinent, Pangaea, formed about 300 million years ago. The previous supercontinent, Pannotia, formed about 600 million years ago, and its dispersal formed the fragments that ultimately collided to form Pangaea. But beyond this the time span between supercontinents becomes more irregular. For example, the supercontinent before Pannotia, Rodinia, existed ~1.1 billion to ~750 million years ago - a mere 150 million years before Pannotia. The supercontinent before this was Columbia: ~1.8 to 1.5 billion years ago.^[1][2] And before this was Kenorland: ~2.7 to ~2.1 billion years ago. The first supercontinents were Ur (existed ~3 billion years ago) and Vaalbara (~3.6 to ~2.8 billion years ago).

Analysis of the composition of mineral inclusions inside ancient diamonds suggests that the cycle of supercontinental formation and breakup began roughly 3.0 billion years ago. Before 3.2 billion years ago only diamonds with peridotitic compositions (commonly found in the Earth's mantle) formed whereas after 3.0 billion years ago eclogitic diamonds (rocks from the Earth's surface crust) became prevalent. This is thought to be due to the introduction of eclogite into subcontinental diamond-forming fluids via subduction and continental collision.^[3]

The hypothetical supercontinent cycle is, in some ways, the complement to the Wilson cycle. The latter is named after plate tectonics pioneer J. Tuzo Wilson and describes the periodic opening and closing of ocean basins. Because the oldest seafloor is only 170 million years old, whereas the oldest bit of continental crust goes back to 4 billion years or more, it makes sense to emphasize the much longer record of the planetary pulse that is recorded in the continents.

Effects on sea level

It is known that sea level is generally low when the continents are together and high when they are apart. For example, sea level was low at the time of formation of Pangaea (Permian) and Pannotia (latest Neoproterozoic), and rose rapidly to maxima during Ordovician and Cretaceous times, when the continents were dispersed. This is because the age of the oceanic lithosphere provides a major control on the depth of the ocean basins, and therefore on global sea level. Oceanic lithosphere forms at mid-ocean ridges and moves outwards. As this happens, it conductively cools and shrinks. This cooling and shrinking decreases the thickness and increases the density of the oceanic lithosphere, and the result is the general lowering in elevation of the seafloor away from mid-ocean ridges.^[4] For oceanic lithosphere that is less than about 75 million years old, a simple cooling half-space model of conductive cooling works,^[4] in which the depth of the ocean basins d in areas in which there is no nearby subduction is a simple function of the age of the oceanic lithosphere t. In general,

$d(t) = 2 \sqrt{\kappa t} + h_r$

where κ is the thermal conductivity of the mantle lithosphere, and is approximately 10⁻⁶ [m²/s, and h_r is the depth of the ridge below the ocean surface. After plugging in rough numbers for the sea floor, the equation becomes:

$d(t) = 350 \sqrt{t} + 2500$

where d is in meters and t is in millions of years, so that just-formed crust at the mid-ocean ridges lies at about 2,500 m depth, whereas 50 million-year-old seafloor lies at a depth of about 5000 m.

As the mean level of the sea floor decreases, the volume of the ocean basins increases, and if other factors that can control sea level remain constant, sea level falls. The converse is also true: younger oceanic lithosphere leads to shallower oceans and higher sea levels if other factors remain constant.

Area A can change when continents rift (stretching the continents decreases A and raises sea level) or as a result of continental collision (compressing the continents leads to an increase A and lowers sea level). Increasing sea level will flood the continents, while decreasing sea level will expose continental shelves.

Because the continental shelf has a very low slope, a small increase in sea level will result in a large change in the percent of continents flooded.

If the world ocean on average is young, the seafloor will be relatively shallow, and sea level will be high: more of the continents are flooded. If the world ocean is on average old, seafloor will be relatively deep, and sea level will be low: more of the continents will be exposed.

There is thus a relatively simple relationship between the Supercontinent Cycle and the mean age of the seafloor.

• Supercontinent = lots of old seafloor = low sea level
• Dispersed continents = lots of young seafloor = high sea level

There will also be a climatic effect of the supercontinent cycle that will amplify this further:

• Supercontinent = continental climate dominant = continental glaciation likely = still lower sea level
• Dispersed continents = maritime climate dominant = continental glaciation unlikely = sea level is not lowered by this mechanism

Relation to global tectonics

There is a progression of tectonic regimes that accompany the supercontinent cycle:

During break-up of the supercontinent, rifting environments dominate. This is followed by passive margin environments, while seafloor spreading continues and the oceans grow. This in turn is followed by the development of collisional environments that become increasingly important with time. First collisions are between continents and island arcs, but lead ultimately to continent-continent collisions. This is the situation that was observed during the Paleozoic Supercontinent Cycle and is being observed for the Mesozoic-Cenozoic Supercontinent Cycle, still in progress.

Relation to climate

Main article: Greenhouse and icehouse Earth

There are two types of global earth climates: icehouse and greenhouse. Icehouse is characterized by frequent continental glaciations and severe desert environments. Greenhouse is characterized by warm climates. Both reflect the supercontinent cycle. We are now in a little greenhouse phase of an ice house world.^[5]

• Icehouse climate
  • Continents moving together
  • Sea level low due to lack of seafloor production
  • Climate cooler, arid
  • Associated with Aragonite seas
  • Formation of Supercontinents

• Greenhouse climate
  • Continents dispersed
  • Sea level high
  • High level of sea floor spreading
  • Relatively large amounts of CO₂ production at oceanic rifting zones
  • Climate warm and humid
  • Associated with calcite seas

Periods of icehouse climate: much of Neoproterozoic, late Paleozoic, late Cenozoic.

Periods of greenhouse climate: Early Paleozoic, Mesozoic-early Cenozoic.

Relation to evolution

The principal mechanism for evolution is natural selection among diverse populations. As genetic drift occurs more frequently in small populations, diversity is an observed consequence of isolation. Less isolation, and thus less diversification, occurs when the continents are all together, producing both one continent and one ocean with one coast. In Latest Neoproterozoic to Early Paleozoic times, when the tremendous proliferation of diverse metazoa occurred, isolation of marine environments resulted from the breakup of Pannotia.

An arrangement of N-S continents and oceans leads to much more diversity and isolation than E-W oceans and continents. This forms zones that are separated by water or land and that merge into climatically different zones along communication routes to the north and south. Formation of similar tracts of continents and ocean basins, only oriented E-W would lead to much less isolation, diversification, and slower evolution. Through the Cenozoic, isolation has been maximized by an arrangement of N-S ocean basins and continents.

Diversity, as measured by the number of families, follows the supercontinent cycle very well.

References

1. ^ Zhao, Guochun; Cawood, Peter A.; Wilde, Simon A.; Sun, M. (2002). "Review of global 2.1–1.8 Ga orogens: implications for a pre-Rodinia supercontinent". Earth-Science Reviews 59 (1-4): 125–162. Bibcode 2002ESRv...59..125Z. doi:10.1016/S0012-8252(02)00073-9. 
2. ^ Zhao, Guochun; Sun, M.; Wilde, Simon A.; Li, S.Z. (2004). "A Paleo-Mesoproterozoic supercontinent: assembly, growth and breakup". Earth-Science Reviews 67 (1-2): 91–123. doi:10.1016/j.earscirev.2004.02.003. 
3. ^ Shirey, S. B.; Richardson, S. H. (2011). "Start of the Wilson Cycle at 3 Ga Shown by Diamonds from Subcontinental Mantle". Science 333 (6041): 434–436. doi:10.1126/science.1206275. PMID 21778395.  edit
4. ^ ^{a b} Parsons, Barry; Sclater, John G.. "An analysis of the variation of ocean floor bathymetry and heat flow with age". Journal of Geophysical Research (American Geophysical Union) 82 (B5): 802–827. Bibcode 1977JGR....82..802P. 
5. ^ Read, J.Fred (2001). "Record of ancient climates can be a map to riches". Science From Virginia Tech. http://www.research.vt.edu/resmag/sciencecol/2001old_weather.html. Retrieved 2011-05-04. 

• Gurnis, M. (1988). "Large-scale mantle convection and the aggregation and dispersal of supercontinents". Nature 332 (6166): 695–699. Bibcode 1988Natur.332..695G. doi:10.1038/332695a0. 
• Murphy, J. B.; Nance, R. D. (1992). "Supercontinents and the origin of mountain belts". Scientific American 266 (4): 84–91. 
• Nance, R. D.; Worsley, T. R.; Moody, J. B. (1988). "The supercontinent cycle". Scientific American 259 (1): 72–79. 

External links

• Reconstructions from "Paleomap Project"
• Plate reconstruction movies from UTIG 'PLATES' project
• A Tectonic Rock Cycle